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364

Ocean Acidification and Other Ocean Changes

13

Climate Science Special ReportU.S. Global Change Research Program

Recommended Citation for ChapterJewett, L. and A. Romanou, 2017: Ocean acidification and other ocean changes. In: Climate Science Spe-cial Report: Fourth National Climate Assessment, Volume I [Wuebbles, D.J., D.W. Fahey, K.A. Hibbard, D.J. Dokken, B.C. Stewart, and T.K. Maycock (eds.)]. U.S. Global Change Research Program, Washington, DC, USA, pp. 364-392, doi: 10.7930/J0QV3JQB.

KEY FINDINGS1. The world’s oceans have absorbed about 93% of the excess heat caused by greenhouse gas warming

since the mid-20th century, making them warmer and altering global and regional climate feedbacks. Ocean heat content has increased at all depths since the 1960s and surface waters have warmed by about 1.3° ± 0.1°F (0.7° ± 0.08°C) per century globally since 1900 to 2016. Under a higher scenario, a global increase in average sea surface temperature of 4.9° ± 1.3°F (2.7° ± 0.7°C) by 2100 is projected, with even higher changes in some U.S. coastal regions. (Very high confidence)

2. The potential slowing of the Atlantic meridional overturning circulation (AMOC; of which the Gulf Stream is one component)—as a result of increasing ocean heat content and freshwater driven buoyancy changes—could have dramatic climate feedbacks as the ocean absorbs less heat and CO2 from the atmosphere. This slowing would also affect the climates of North America and Europe. Any slowing documented to date cannot be directly tied to anthropogenic forcing primarily due to lack of adequate observational data and to challenges in modeling ocean circulation changes. Under a higher scenario (RCP8.5) in CMIP5 simulations, the AMOC weakens over the 21st century by 12% to 54% (low confidence).

3. The world’s oceans are currently absorbing more than a quarter of the CO2 emitted to the atmo-sphere annually from human activities, making them more acidic (very high confidence), with potential detrimental impacts to marine ecosystems. In particular, higher-latitude systems typically have a lower buffering capacity against pH change, exhibiting seasonally corrosive conditions sooner than low-latitude systems. Acidification is regionally increasing along U.S. coastal systems as a result of upwelling (for example, in the Pacific Northwest) (high confidence), changes in freshwater inputs (for example, in the Gulf of Maine) (medium confidence), and nutrient input (for example, in agricultural watersheds and urbanized estuaries) (high confidence). The rate of acidification is unparalleled in at least the past 66 million years (medium confidence). Under the higher scenario (RCP8.5), the global average surface ocean acidity is projected to increase by 100% to 150% (high confidence).

4. Increasing sea surface temperatures, rising sea levels, and changing patterns of precipitation, winds, nutrients, and ocean circulation are contributing to overall declining oxygen concentrations at inter-mediate depths in various ocean locations and in many coastal areas. Over the last half century, major oxygen losses have occurred in inland seas, estuaries, and in the coastal and open ocean (high confi-dence). Ocean oxygen levels are projected to decrease by as much as 3.5% under the higher scenario (RCP8.5) by 2100 relative to preindustrial values (high confidence).

http://doi.org/10.7930/J0QV3JQB

13 | Ocean Acidification and Other Ocean Changes

365 Climate Science Special ReportU.S. Global Change Research Program

13.0 A Changing Ocean Anthropogenic perturbations to the global Earth system have included important alter-ations in the chemical composition, tempera-ture, and circulation of the oceans. Some of these changes will be distinguishable from the background natural variability in nearly half of the global open ocean within a decade, with important consequences for marine ecosys-tems and their services.1 However, the time-frame for detection will vary depending on the parameter featured.2, 3

13.1 Ocean Warming

13.1.1 General BackgroundApproximately 93% of excess heat energy trapped since the 1970s has been absorbed into the oceans, lessening atmospheric warm-ing and leading to a variety of changes in ocean conditions, including sea level rise and ocean circulation (see Ch. 2: Physical Driv-ers of Climate Change, Ch. 6: Temperature Change, and Ch. 12: Sea Level Rise in this report).1, 4 This is the result of the high heat ca-pacity of seawater relative to the atmosphere, the relative area of the ocean compared to the land, and the ocean circulation that enables the transport of heat into deep waters. This large heat absorption by the oceans moderates the effects of increased anthropogenic green-house emissions on terrestrial climates while altering the fundamental physical properties of the ocean and indirectly impacting chem-ical properties such as the biological pump through increased stratification.1, 5 Although upper ocean temperature varies over short- and medium timescales (for example, seasonal and regional patterns), there are clear long-term increases in surface temperature and ocean heat content over the past 65 years.4, 6, 7

13.1.2 Ocean Heat ContentOcean heat content (OHC) is an ideal variable to monitor changing climate as it is calculat-ed using the entire water column, so ocean

warming can be documented and compared between particular regions, ocean basins, and depths. However, for years prior to the 1970s, estimates of ocean uptake are confined to the upper ocean (up to 700 m) due to sparse spa-tial and temporal coverage and limited ver-tical capabilities of many of the instruments in use. OHC estimates are improved for time periods after 1970 with increased sampling coverage and depth.4, 8 Estimates of OHC have been calculated going back to the 1950s us-ing averages over longer time intervals (i.e., decadal or 5-year intervals) to compensate for sparse data distributions, allowing for clear long-term trends to emerge (e.g., Levitus et al. 20127).

From 1960 to 2015, OHC significantly in-creased for both 0–700 and 700–2,000 m depths, for a total ocean warming of about 33.5 ± 7.0 × 1022 J (a net heating of 0.37 ± 0.08 W/m2; Figure 13.1).6 During this period, there is evidence of an acceleration of ocean warming beginning in 1998,9 with a total heat increase of about 15.2 × 1022 J.6 Robust ocean warming occurs in the upper 700 m and is slow to penetrate into the deep ocean. However, the 700–2,000 m depths constitute an increasing portion of the total ocean ener-gy budget as compared to the surface ocean (Figure 13.1).6 The role of the deep ocean (below 2,000 m [6,600 ft]) in ocean heat uptake remains uncertain, both in the magnitude but also the sign of the uptake.10, 11 Penetration of surface waters to the deep ocean is a slow process, which means that while it takes only about a decade for near-surface temperatures to respond to increased heat energy, the deep ocean will continue to warm, and as a result sea levels will rise for centuries to millennia even if all further emissions cease.4



Several sources have documented warming in all ocean basins from 0–2,000 m depths over the past 50 years (Figure 13.2).6, 7, 12 Annual fluctu-ations in surface temperatures and OHC are attributed to the combination of a long-term secular trend and decadal and smaller time scale variations, such as the Pacific Decadal Os-cillation (PDO) and the Atlantic Multidecadal Oscillation (AMO) (Ch. 5: Circulation & Vari-ability; Ch. 12: Sea Level Rise).13, 14 The trans-port of heat to the deep ocean is likely linked to the strength of the Atlantic Meridional Over-turning Circulation (see Section 13.2.1), where the Atlantic and Southern Ocean accounts for the dominant portion of total OHC change at the 700–2,000 m depth.6, 8, 9, 15 Decadal variabil-

ity in ocean heat uptake is mostly attributed to ENSO phases (with El Niños warming and La Niñas cooling). For instance, La Niña con-ditions over the past decade have led to colder ocean temperatures in the eastern tropical Pacific.6, 8, 9, 16 For the Pacific and Indian Oceans, the decadal shifts are primarily observed in the upper 350 m depth, likely due to shallow subtropical circulation, leading to an abrupt increase of OHC in the Indian Ocean carried by the Indonesian throughflow from the Pacific Ocean over the last decade.9 Although there is natural variability in ocean temperature, there remain clear increasing trends due to anthropo-genic influences.

Figure 13.1: Global Ocean heat content change time series. Ocean heat content from 0 to 700 m (blue), 700 to 2,000 m (red), and 0 to 2,000 m (dark gray) from 1955 to 2015 with an uncertainty interval of ±2 standard deviations shown in shading. All time series of the analysis performed by Cheng et al.6 are smoothed by a 12-month running mean filter, relative to the 1997–2005 base period. (Figure source: Cheng et al. 20176).

1950 1960 1970 1980 1990 2000 2010−35

−30

−25

−20

−15

−10

−5

0

5

10

15

Year

Oce

an h

eat c

onte

nt (1

022 J

oule

s)

700−2000m0−700m

0−2000m



13.1.3 Sea Surface Temperature and U.S. Regional WarmingIn addition to OHC, sea surface temperature (SST) measurements are widely available. SST measurements are useful because 1) the measurements have been taken over 150 years (albeit using different platforms, instruments, and depths through time); 2) SST reflects the lower boundary condition of the atmosphere; and 3) SST can be used to predict specific regional impacts of global warming on ter-restrial and coastal systems.15, 17, 18 Globally, surface ocean temperatures have increased by 1.3° ± 0.1°F (0.70° ± 0.08°C) per century from 1900 to 2016 for the Extended Reconstructed Sea Surface Temperature version 4 (ERSST v4) record.19 All U.S. coastal waters have warmed by more than 0.7°F (0.4°C) over this period as shown in both Table 13.1 and Chapter 6: Temperature Change, Figure 6.6. During the past 60 years, the rates of increase of SSTs for the coastal waters of three U.S. regions were above the global average rate. These includ-ed the waters around Alaska, the Northeast, and the Southwest (Table 13.1). Over the last decade, some regions have experienced

increased high ocean temperature anomalies. SST in the Northeast has warmed faster than 99% of the global ocean since 2004, and a peak temperature for the region in 2012 was part of a large “ocean heat wave” in the Northwest Atlantic that persisted for nearly 18 months.20,

21 Projections indicate that the Northeast will continue to warm more quickly than other ocean regions through the end of the cen-tury.22 In the Northwest, a resilient ridge of high pressure over the North American West Coast suppressed storm activity and mixing, which intensified heat in the upper ocean in a phenomenon known as “The Blob”.23 Anom-alously warm waters persisted in the coastal waters of the Alaskan and Pacific Northwest from 2013 until 2015. Under a higher scenario (RCP8.5), SSTs are projected to increase by an additional 4.9°F (2.7°C) by 2100 (Figure 13.3), whereas for a lower scenario (RCP4.5) the SST increase would be 2.3°F (1.3°C).24 In all U.S. coastal regions, the warming since 1901 is detectable compared to natural variability and attributable to anthropogenic forcing, accord-ing to an analysis of the CMIP5 models (Ch. 6: Temperature Change, Figure 6.5).

Figure 13.2: Ocean heat content changes from 1960 to 2015 for different ocean basins for 0 to 2,000 m depths. Time series is relative to the 1997–1999 base period and smoothed by a 12-month running filter by Cheng et al.6 The curves are additive, and the ocean heat content changes in different ocean basins are shaded in different colors (Figure source: Cheng et al. 20176).

OH

C (0

–200

0 m

) 1022

J

−15

1960

Indian OceanNorth AtlanticTropical/subtropical AtlanticNorth PacificTropical/subtropical PacificSouthern oceans

19801900 1990 2000 2010

−10

−5

0

5

10

15

20



Table 13.1. Historical sea surface temperature trends (°C per century) and projected trends by 2080 (°C) for eight U.S. coastal regions and globally. Historical temperature trends are presented for the 1900–2016 and 1950–2016 periods with 95% confidence level, observed using the Extended Reconstructed Sea Surface Temperature version 4 (ERSSTv4).19 Global and regional predictions are calculated for lower and higher scenarios (RCP4.5 and RCP8.5, respectively) with 80% spread of all the CMIP5 members compared to the 1976–2005 period.151 The historical trends were analyzed for the latitude and longi-tude in the table, while the projected trends were analyzed for the California Current in-stead of the Northwest and Southwest separately and for the Bering Sea in Alaska (NOAA).

Region Latitude and Longitude

Historical Trend(°C/100 years)

Projected Trend 2080 (relative to

1976–2005 climate) (°C)

1900–2016 1950–2016 RCP4.5 RCP8.5

Global 0.70 ± 0.08 1.00 ± 0.11 1.3 ± 0.6 2.7 ± 0.7

Alaska 50°–66°N, 150°–170°W 0.82 ± 0.26 1.22 ± 0.59 2.5 ± 0.6 3.7 ± 1.0

Northwest (NW)

40°–50°N, 120°–132°W 0.64 ± 0.30 0.68 ± 0.70

1.7 ± 0.4 2.8 ± 0.6Southwest

(SW)30°–40°N,

116°–126°W 0.73 ± 0.33 1.02 ± 0.79

Hawaii (HI) 18°–24°N, 152°–162°W 0.58 ± 0.19 0.46 ± 0.39 1.6 ± 0.4 2.8 ± 0.6

Northeast (NE) 36°–46°N, 64°–76°W 0.63 ± 0.31 1.10 ± 0.71 2.0 ± 0.3 3.2 ± 0.6

Southeast (SE) 24°–34°N, 64°–80°W 0.40 ± 0.18 0.13 ± 0.34 1.6 ± 0.3 2.7 ± 0.4

Gulf of Mexico (GOM)

20°–30°N, 80°–96°W 0.52 ± 0.14 0.37 ± 0.27 1.6 ± 0.3 2.8 ± 0.3

Caribbean 10°–20°N, 66°–86°W 0.76 ± 0.15 0.77 ± 0.32 1.5 ± 0.4 2.6 ± 0.3



13.1.4 Ocean Heat FeedbackThe residual heat not taken up by the oceans increases land surface temperatures (approx-imately 3%) and atmospheric temperatures (approximately 1%), and melts both land and sea ice (approximately 3%), leading to sea level rise (see Ch. 12: Sea Level Rise).4, 6, 25 The meltwater from land and sea ice amplifies further subsurface ocean warming and ice shelf melting, primarily due to increased ther-mal stratification, which reduces the ocean’s efficiency in transporting heat to deep waters.4 Surface ocean stratification has increased by about 4% during the period 1971–201026 due to thermal heating and freshening from increased freshwater inputs (precipitation and evaporation changes and land and sea ice melting). The increase of ocean stratification will contribute to further feedback of ocean warming and, indirectly, mean sea level. In addition, increases in stratification are associ-ated with suppression of tropical cyclone in-tensification,27 retreat of the polar ice sheets,28 and reductions of the convective mixing at higher latitudes that transports heat to the

deep ocean through the Atlantic Meridional Overturning Circulation.29 Ocean heat uptake therefore represents an important feedback that will have a significant influence on future shifts in climate (see Ch. 2: Physical Drivers of Climate Change).

13.2 Ocean Circulation

13.2.1 Atlantic Meridional Overturning CirculationThe Atlantic Meridional Overturning Circula-tion (AMOC) refers to the three-dimensional, time-dependent circulation of the Atlantic Ocean, which has been a high priority top-ic of study in recent decades. The AMOC plays an important role in climate through its transport of heat, freshwater, and carbon (e.g., Johns et al. 2011;30 McDonagh et al. 2015;31 Talley et al. 201632). AMOC-associated poleward heat transport substantially con-tributes to North American and continental European climate (see Ch. 5: Circulation and Variability). The Gulf Stream, in contrast to other western boundary currents, is expected to slow down because of the weakening of the AMOC, which would impact the Euro-

Figure 13.3: Projected changes in sea surface temperature (°C) for the coastal United States under the higher scenar-io (RCP8.5). Projected anomalies for the 2050–2099 period are calculated using a comparison from the average sea surface temperatures over 1956–2005. Projected changes are examined using the Coupled Model Intercomparison Project Phase 5 (CMIP5) suite of model simulations. (Figure source: NOAA).

CMIP5 ENSMN RCP8.5 Anomaly

0.0 0.3 0.6 0.9 1.2 1.5

°C

1.8 2.1 2.4 2.7 3.0 3.3

(2050–2099)−(1956–2005)CMIP5 ENSMN RCP8.5 Anomaly

(2050–2099)–(1956–2005)



pean climate.33 Variability in the AMOC has been attributed to wind forcing on intra-an-nual time scales and to geostrophic forces on interannual to decadal timescales.34 Increased freshwater fluxes from melting Arctic Sea and land ice can weaken open ocean convection and deep-water formation in the Labrador and Irminger Seas, which could weaken the AMOC (Ch. 11: Arctic Changes; also see Ch. 5, Section 5.2.3: North Atlantic Oscillation and Northern Annular Mode).29, 33

While one recent study has suggested that the AMOC has slowed since preindustrial times29 and another suggested slowing on faster time scales,35 there is at present insufficient obser-vational evidence to support a finding of long term slowdown of AMOC strength over the 20th century4 or within the last 50 years34 as decadal ocean variability can obscure long-term trends. Some studies show long-term trends,36, 37 but the combination of sparse data and large seasonal variability may also lead to incorrect interpretations (e.g., Kanzow et al. 201038). Several recent high resolution modeling studies constrained with the limit-ed existing observational data39 and/or with reconstructed freshwater fluxes40 suggest that the recently observed AMOC slowdown at 26°N (off the Florida coast) since 2004 (e.g., as described in Smeed et al. 201435) is mainly due to natural variability, and that anthropogenic forcing has not yet caused a significant AMOC slowdown. In addition, direct observations of the AMOC in the South Atlantic fail to unam-biguously demonstrate anthropogenic trends (e.g., Dong et al. 2015;41 Garzoli et al. 201342).

Under a higher scenario (RCP8.5) in CMIP5 simulations, it is very likely that the AMOC will weaken over the 21st century. The project-ed decline ranges from 12% to 54%,43 with the range width reflecting substantial uncertainty in quantitative projections of AMOC behavior. In lower scenarios (like RCP4.5), CMIP5 mod-

els predict a 20% weakening of the AMOC during the first half of the 21st century and a stabilization and slight recovery after that.44 The projected slowdown of the AMOC will be counteracted by the warming of the deep ocean (below 700 m [2,300 ft]), which will tend to strengthen the AMOC.45 The situa-tion is further complicated due to the known bias in coupled climate models related to the direction of the salinity transport in models versus observations, which is an indicator of AMOC stability (e.g., Drijhout et al. 2011;46 Bryden et al. 2011;47 Garzoli et al. 201342). Some argue that coupled climate models should be corrected for this known bias and that AMOC variations could be even larger than the gradual decrease most models predict if the AMOC were to shut down completely and “flip states”.48 Any AMOC slowdown could result in less heat and CO2 absorbed by the ocean from the atmosphere, which is a posi-tive feedback to climate change (also see Ch. 2: Physical Drivers of Climate Change).49, 50, 51

13.2.2 Changes in Salinity StructureAs a response to warming, increased atmo-spheric moisture leads to stronger evapora-tion or precipitation in terrestrial and oceanic environments and melting of land and sea ice. Approximately 80% of precipitation/evapo-ration events occur over the ocean, leading to patterns of higher salt content or freshwater anomalies and changes in ocean circulation (see Ch. 2: Physical Drivers of Climate Change and Ch. 6: Temperature Change).52 Over 1950–2010, average global amplification of the sur-face salinity pattern amounted to 5.3%; where fresh regions in the ocean became fresher and salty regions became saltier.53 However, the long-term trends of these physical and chemi-cal changes to the ocean are difficult to isolate from natural large-scale variability. In partic-ular, ENSO displays particular salinity and precipitation/evaporation patterns that skew the trends. More research and data are neces-



sary to better model changes to ocean salinity. Several models have shown a similar spatial structure of surface salinity changes, including general salinity increases in the subtropical gyres, a strong basin-wide salinity increase in the Atlantic Ocean, and reduced salinity in the western Pacific warm pools and the North Pacific subpolar regions.52, 53 There is also a stronger distinction between the upper salty thermocline and fresh intermediate depth through the century. The regional changes in salinity to ocean basins will have an overall impact on ocean circulation and net primary production, leading to corresponding carbon export (see Ch. 2: Physical Drivers of Climate Change). In particular, the freshening of the Arctic Ocean due to melting of land and sea ice can lead to buoyancy changes which could slow down the AMOC (see Section 13.2.1).

13.2.3 Changes in UpwellingSignificant changes to ocean stratification and circulation can also be observed regionally, along the eastern ocean boundaries and at the equator. In these areas, wind-driven upwell-ing brings colder, nutrient- and carbon-rich water to the surface; this upwelled water is more efficient in heat and anthropogenic CO2 uptake. There is some evidence that coastal upwelling in mid- to high-latitude eastern boundary regions has increased in intensity and/or frequency,54 but in more tropical areas of the western Atlantic, such as in the Carib-bean Sea, it has decreased between 1990 and 2010.55, 56 This has led to a decrease in primary productivity in the southern Caribbean Sea.55 Within the continental United States, the Cali-fornia Current is experiencing fewer (by about 23%–40%) but stronger upwelling events.57, 58,

59 Stronger offshore upwelling combined with cross-shelf advection brings nutrients from the deeper ocean but also increased offshore transport.60 The net nutrient load in the coastal regions is responsible for increased productiv-ity and ecosystem function.

IPCC 2013 concluded that there is low con-fidence in the current understanding of how eastern upwelling systems will be altered under future climate change because of the ob-scuring role of multidecadal climate variabil-ity.26 However, subsequent studies show that by 2100, upwelling is predicted to start earlier in the year, end later, and intensify in three of the four major eastern boundary upwelling systems (not in the California Current).61 In the California Current, upwelling is projected to intensify in spring but weaken in summer, with changes emerging from the envelope of natural variability primarily in the second half of the 21st century.62 Southern Ocean upwell-ing will intensify while the Atlantic equato-rial upwelling systems will weaken.57, 61 The intensification is attributed to the strength-ening of regional coastal winds as observa-tions already show,58 and model projections under the higher scenario (RCP8.5) estimate wind intensifying near poleward boundaries (including northern California Current) and weakening near equatorward boundaries (in-cluding southern California Current) for the 21st century.61, 63

13.3 Ocean Acidification

13.3.1 General BackgroundIn addition to causing changes in climate, increasing atmospheric levels of carbon di-oxide (CO2) from the burning of fossil fuels and other human activities, including chang-es in land use, have a direct effect on ocean carbonate chemistry that is termed ocean acidification.64, 65 Surface ocean waters absorb part of the increasing CO2 in the atmosphere, which causes a variety of chemical changes in seawater: an increase in the partial pressure of CO2 (pCO2,sw), dissolved inorganic carbon (DIC), and the concentration of hydrogen and bicarbonate ions and a decrease in the concen-tration of carbonate ions (Figure 13.4). In brief, CO2 is an acid gas that combines with water to form carbonic acid, which then dissociates



to hydrogen and bicarbonate ions. Increasing concentrations of seawater hydrogen ions result in a decrease of carbonate ions through their conversion to bicarbonate ions. The con-centration of carbonate ions in seawater affects saturation states for calcium carbonate com-pounds, which many marine species use to build their shells and skeletons. Ocean acidity refers to the concentration of hydrogen ions in ocean seawater regardless of ocean pH, which is fundamentally basic (e.g., pH > 7). Ocean

surface waters have become 30% more acidic over the last 150 years as they have absorbed large amounts of CO2 from the atmosphere,66 and anthropogenically sourced CO2 is gradu-ally invading into oceanic deep waters. Since the preindustrial period, the oceans have ab-sorbed approximately 29% of all CO2 emitted to the atmosphere.67 Oceans currently absorb about 26% of the human-caused CO2 anthro-pogenically emitted into the atmosphere.67

Figure 13.4: Trends in surface (< 50 m) ocean carbonate chemistry calculated from observations obtained at the Hawai‘i Ocean Time-series (HOT) Program in the North Pacific over 1988–2015. The upper panel shows the linked increase in atmospheric (red points) and seawater (blue points) CO2 concentrations. The bottom panel shows a de-cline in seawater pH (black points, primary y-axis) and carbonate ion concentration (green points, secondary y-axis). Ocean chemistry data were obtained from the Hawai‘i Ocean Time-series Data Organization & Graphical System (HOT-DOGS, http://hahana.soest.hawaii.edu/hot/hot-dogs/index.html). (Figure source: NOAA).

Carbon Dioxide

Air

Sea

*calculated from discrete water samples

270

CO32−

(µmo

l/kg)260

250240230220210

19807.807.857.907.958.008.058.108.15300325350375400

pHpp

m

1985 1990 1995 2000Year

2005 2010 2015



13.3.2 Open Ocean AcidificationSurface waters in the open ocean experience changes in carbonate chemistry reflective of large-scale physical oceanic processes (see Ch. 2: Physical Drivers of Climate Change). These processes include both the global uptake of atmospheric CO2 and the shoaling of natural-ly acidified subsurface waters due to vertical mixing and upwelling. In general, the rate of ocean acidification in open ocean surface waters at a decadal time-scale closely approx-imates the rate of atmospheric CO2 increase.68 Large, multidecadal phenomena such as the Atlantic Multidecadal Oscillation and Pacific Decadal Oscillation can add variability to the observed rate of change.68

13.3.3 Coastal AcidificationCoastal shelf and nearshore waters are influ-enced by the same processes as open ocean surface waters such as absorption of atmo-spheric CO2 and upwelling, as well as a num-ber of additional, local-level processes, includ-ing freshwater, nutrient, sulfur, and nitrogen inputs.69, 70 Coastal acidification generally ex-hibits higher-frequency variability and short-term episodic events relative to open-ocean acidification.71, 72, 73, 74 Upwelling is of particular importance in coastal waters, especially along the U.S. West Coast. Deep waters that shoal with upwelling are enriched in CO2 due to up-take of anthropogenic atmospheric CO2 when last in contact with the atmosphere, coupled with deep water respiration processes and lack of gas exchange with the atmosphere.65,

75 Freshwater inputs to coastal waters change seawater chemistry in ways that make it more susceptible to acidification, largely by fresh-ening ocean waters and contributing varying amounts of dissolved inorganic carbon (DIC), total alkalinity (TA), dissolved and particulate organic carbon, and nutrients from riverine and estuarine sources. Coastal waters of the East Coast and mid-Atlantic are far more in-fluenced by freshwater inputs than are Pacific

Coast waters.76 Coastal waters can episodically experience riverine and glacial melt plumes that create conditions in which seawater can dissolve calcium carbonate structures.77, 78 While these processes have persisted histor-ically, climate-induced increases in glacial melt and high-intensity precipitation events can yield larger freshwater plumes than have occurred in the past. Nutrient runoff can increase coastal acidification by creating conditions that enhance biological respiration. In brief, nutrient loading typically promotes phytoplankton blooms, which, when they die, are consumed by bacteria. Bacteria respire CO2 and thus bacterial blooms can result in acidification events whose intensity depends on local hydrographic conditions, including water column stratification and residence time.72 Long-term changes in nutrient loading, precipitation, and/or ice melt may also impart long-term, secular changes in the magnitude of coastal acidification.

13.3.4 Latitudinal VariationOcean carbon chemistry is highly influ-enced by water temperature, largely because the solubility of CO2 in seawater increases as water temperature declines. Thus, cold, high-latitude surface waters can retain more CO2 than warm, lower-latitude surface wa-ters.76, 79 Because carbonate minerals also more readily dissolve in colder waters, these waters can more regularly become undersaturated with respect to calcium carbonate whereby mineral dissolution is energetically favored. This chemical state, often referred to as sea-water being “corrosive” to calcium carbonate, is important when considering the ecological implications of ocean acidification as many species make structures such as shells and skeletons from calcium carbonate. Seawater conditions undersaturated with respect to calcium carbonate are common at depth, but currently and historically rare at the surface and near-surface.80 Some high-latitude surface



and near-surface waters now experience such corrosive conditions, which are rarely docu-mented in low-latitude surface or near-surface systems. For example, corrosive conditions at a range of ocean depths have been docu-mented in the Arctic and northeastern Pacific Oceans.74, 79, 81, 82 Storm-induced upwelling could cause undersaturation in tropical ar-eas in the future.83 It is important to note that low-latitude waters are experiencing a greater absolute rate of change in calcium carbonate saturation state than higher latitudes, though these low-latitude waters are not approaching the undersaturated state except within near-shore or some benthic habitats.84

13.3.5 Paleo EvidenceEvidence suggests that the current rate of ocean acidification is the fastest in the last 66 million years (the K-Pg boundary) and pos-sibly even the last 300 million years (when the first pelagic calcifiers evolved providing proxy information and also a strong carbonate buffer, characteristic of the modern ocean).85, 86 The Paleo-Eocene Thermal Maximum (PETM; around 56 million years ago) is often refer-enced as the closest analogue to the present, although the overall rate of change in CO2 conditions during that event (estimated be-tween 0.6 and 1.1 GtC/year) was much lower than the current increase in atmospheric CO2 of 10 GtC/year.86, 87 The relatively slower rate of atmospheric CO2 increase at the PETM like-ly led to relatively small changes in carbonate ion concentration in seawater compared with the contemporary acidification rate, due to the ability of rock weathering to buffer the change over the longer time period.86 Some of the pre-sumed acidification events in Earth’s history have been linked to selective extinction events suggestive of how guilds of species may re-spond to the current acidification event.85

13.3.6 Projected ChangesProjections indicate that by the end of the century under higher scenarios, such as SRES A1FI or RCP8.5, open-ocean surface pH will decline from the current average level of 8.1 to a possible average of 7.8 (Figure 13.5).1 When the entire ocean volume is considered under the same scenario, the volume of waters undersaturated with respect to calcium car-bonate could expand from 76% in the 1990s to 91% in 2100, resulting in a shallowing of the saturation horizons—depths below which un-dersaturation occurs.1, 88 Saturation horizons, which naturally vary among ocean basins, influence ocean carbon cycles and organisms with calcium carbonate structures, especially as they shoal into the zones where most biota lives.81, 89 As discussed above, for a variety of reasons, not all ocean and coastal regions will experience acidification in the same way depending on other compounding factors. For instance, recent observational data from the Arctic Basin show that the Beaufort Sea became undersaturated, for part of the year, with respect to aragonite in 2001, while other continental shelf seas in the Arctic Basin are projected to do so closer to the middle of the century (e.g., the Chukchi Sea in about 2033 and Bering Sea in about 2062).90 Deviation from the global average rate of acidification will be especially true in coastal and estuarine areas where the rate of acidification is influ-enced by other drivers than atmospheric CO2, some of which are under the control of local management decisions (for example, nutrient pollution loads).



13.4 Ocean Deoxygenation

13.4.1 General BackgroundOxygen is essential to most life in the ocean, governing a host of biogeochemical and bi-ological processes. Oxygen influences meta-bolic, physiological, reproductive, behavioral, and ecological processes, ultimately shaping the composition, diversity, abundance, and distribution of organisms from microbes to whales. Increasingly, climate-induced oxygen loss (deoxygenation) associated with ocean warming and reduced ventilation to deep wa-ters has become evident locally, regionally, and globally. Deoxygenation can also be attributed to anthropogenic nutrient input, especially in the coastal regions, where the nutrients can lead to the proliferation of primary production and, consequently, enhanced drawdown of dissolved oxygen by microbes.91 In addition, acidification (Section 13.2) can co-occur with deoxygenation as a result of warming-en-hanced biological respiration.92 As aerobic organisms respire, O2 is consumed and CO2 is produced. Understanding the combined effect of both low O2 and low pH on marine ecosys-tems is an area of active research.93 Warming also raises biological metabolic rates which, in combination with intensified coastal and estuarine stratification, exacerbates eutrophi-cation-induced hypoxia. We now see earlier

onset and longer periods of seasonal hypoxia in many eutrophic sites, most of which occur in areas that are also warming.91

13.4.2 Climate Drivers of Ocean DeoxygenationGlobal ocean deoxygenation is a direct effect of warming. Ocean warming reduces the solubility of oxygen (that is, warmer water can hold less oxygen) and changes physical mix-ing (for example, upwelling and circulation) of oxygen in the oceans. The increased tem-perature of global oceans accounts for about 15% of current global oxygen loss,94 although changes in temperature and oxygen are not uniform throughout the ocean.15 Warming also exerts direct influence on thermal stratification and enhances salinity stratification through ice melt and climate change-associated precip-itation effects. Intensified stratification leads to reduced ventilation (mixing of oxygen into the ocean interior) and accounts for up to 85% of global ocean oxygen loss.94 Effects of ocean temperature change and stratification on oxy-gen loss are strongest in intermediate or mode waters at bathyal depths (in general, 200–3,000 m) and also nearshore and in the open ocean; these changes are especially evident in tropical and subtropical waters globally, in the Eastern Pacific,95 and in the Southern Ocean.94

Figure 13.5: Predicted change in sea surface pH in 2090–2099 relative to 1990–1999 under the higher scenario (RCP8.5), based on the Community Earth System Models–Large Ensemble Experiments CMIP5 (Figure source: adapted from Bopp et al. 201324).

Surface pH in 2090s (RCP8.5, changes from 1990s)

−0.5

−0.4

−0.3

−0.2Surface pH in 2090s (RCP8.5, changes from 1990s)



There are also other, less direct effects of global temperature increase. Warming on land reduc-es terrestrial plant water efficiency (through effects on stomata; see Ch. 8: Drought, Floods, and Wildfires, Key Message 3), leading to greater runoff, on average, into coastal zones (see Ch. 8: Drought, Floods, and Wildfires for other hydrological effects of warming) and further enhancing hypoxia potential because greater runoff can mean more nutrient trans-port (See Ch. 2: Physical Drivers of Climate Change).96, 97 Estuaries, especially ones with minimal tidal mixing, are particularly vulner-able to oxygen-depleted dead zones from the enhanced runoff and stratification. Warming can induce dissociation of frozen methane in gas hydrates buried on continental mar-gins, leading to further drawdown of oxygen through aerobic methane oxidation in the water column.98 On eastern ocean boundaries, warming can enhance the land–sea tempera-ture differential, causing increased upwelling due to higher winds with (a) greater nutrient input leading to production, sinking, decay, and biochemical drawdown of oxygen and (b) upwelling of naturally low-oxygen, high-CO2 waters onto the upper slope and shelf environ-ments.58, 65 However, in the California Current, upwelling intensification has occurred only in the poleward regions (north of San Francisco), and the drivers may not be associated with land–sea temperature differences.63 Taken to-gether, the effects of warming are manifested as low-oxygen water in open oceans are being transported to and upwelled along coastal regions. These low-oxygen upwelled waters are then coupled with eutrophication-induced hypoxia, further reducing oxygen content in coastal areas.

Changes in precipitation, winds, circulation, airborne nutrients, and sea level can also contribute to ocean deoxygenation. Project-ed increases in precipitation in some regions will intensify stratification, reducing vertical

mixing and ventilation, and intensify nutrient input to coastal waters through excess runoff, which leads to increased algal biomass and concurrent dissolved oxygen consumption via community respiration.99 Coastal wetlands that might remove these nutrients before they reach the ocean may be lost through rising sea level, further exacerbating hypoxia.97 Some observations of oxygen decline are linked to regional changes in circulation involving low-oxygen water masses. Enhanced fluxes of airborne iron and nitrogen are interacting with natural climate variability and contributing to fertilization, enhanced respiration, and oxygen loss in the tropical Pacific.100

13.4.3 Biogeochemical Feedbacks of Deoxygen-ation to Climate and Elemental CyclesClimate patterns and ocean circulation have a large effect on global nitrogen and oxygen cy-cles, which in turn affect phosphorus and trace metal availability and generate feedbacks to the atmosphere and oceanic production. Global ocean productivity may be affected by climate-driven changes below the tropical and subtropical thermocline which control the volume of suboxic waters (< 5 micromolar O2), and consequently the loss of fixed nitro-gen through denitrification.101, 102 The extent of suboxia in the open ocean also regulates the production of the greenhouse gas nitrous oxide (N2O); as oxygen declines, greater N2O production may intensify global warming, as N2O is about 310 times more effective at trapping heat than CO2 (see Ch. 2: Physical Drivers of Climate Change, Section 2.3.2).103, 104 Production of hydrogen sulfide (H2S, which is highly toxic) and intensified phosphorus recy-cling can occur at low oxygen levels.105 Other feedbacks may emerge as oxygen minimum zone (OMZ) shoaling diminishes the depths of diurnal vertical migrations by fish and inver-tebrates, and as their huge biomass and associ-ated oxygen consumption deplete oxygen.106



13.4.4 Past TrendsOver hundreds of millions of years, oxygen has varied dramatically in the atmosphere and ocean and has been linked to biodiversity gains and losses.107, 108 Variation in oxygen-ation in the paleo record is very sensitive to climate—with clear links to temperature and often CO2 variation.109 OMZs expand and con-tract in synchrony with warming and cooling events, respectively.110 Episodic climate events that involve rapid temperature increases over decades, followed by a cool period lasting a few hundred years, lead to major fluctuations in the intensity of Pacific and Indian Ocean OMZs (i.e., DO of < 20 µM). These events are associated with rapid variations in North At-lantic deep water formation.111 Ocean oxygen fluctuates on glacial-interglacial timescales of thousands of years in the Eastern Pacific.112, 113

13.4.5 Modern Observations (last 50+ years)Long-term oxygen records made over the last 50 years reflect oxygen declines in inland seas,114, 115, 116 in estuaries,117, 118 and in coast-al waters.119, 120, 121, 122 The number of coastal, eutrophication-induced hypoxic sites in the United States has grown dramatically over the past 40 years.123 Over larger scales, global syn-theses show hypoxic waters have expanded by 4.5 million km2 at a depth of 200 m,95 with widespread loss of oxygen in the Southern

Ocean,94 Western Pacific,124 and North Atlan-tic.125 Overall oxygen declines have been great-er in coastal oceans than in the open ocean126 and often greater inshore than offshore.127 The emergence of a deoxygenation signal in regions with naturally high oxygen variability will unfold over longer time periods (20–50 years from now).128

13.4.6 Projected ChangesGlobal ModelsGlobal models generally agree that ocean deoxygenation is occurring; this finding is also reflected in in situ observations from past 50 years. Compilations of 10 Earth Sys-tem models predict a global average loss of oxygen of −3.5% (higher scenario, RCP8.5) to −2.4% (lower scenario, RCP4.5) by 2100, but much stronger losses regionally, and in intermediate and mode waters (Figure 13.6).24 The North Pacific, North Atlantic, Southern Ocean, subtropical South Pacific, and South Indian Oceans all are expected to experience deoxygenation, with O2 decreases of as much as 17% in the North Pacific by 2100 for the RCP8.5 pathway. However, the tropical Atlan-tic and tropical Indian Oceans show increasing O2 concentrations. In the many areas where oxygen is declining, high natural variability makes it difficult to identify anthropogenically forced trends.128



Regional ModelsRegional models are critical because many oxygen drivers are local, influenced by ba-thymetry, winds, circulation, and fresh water and nutrient inputs. Most eastern boundary upwelling areas are predicted to experience intensified upwelling to 2100,61 although on the West Coast projections for increasing upwelling for the northern California Current occur only north of San Francisco (see Section 13.2.3).

Particularly notable for the western United States, variation in trade winds in the eastern Pacific Ocean can affect nutrient inputs, lead-ing to centennial periods of oxygen decline or oxygen increase distinct from global oxygen decline.129 Oxygen dynamics in the Eastern Tropical Pacific are highly sensitive to equato-rial circulation changes.130

Regional modeling also shows that year-to-year variability in precipitation in the central United States affects the nitrate–N flux by the Mississippi River and the extent of hypox-

ia in the Gulf of Mexico.131 A host of climate influences linked to warming and increased precipitation are predicted to lower dissolved oxygen in Chesapeake Bay.132

13.5 Other Coastal Changes

13.5.1 Sea Level RiseSea level is an important variable that affects coastal ecosystems. Global sea level rose rap-idly at the end of the last glaciation, as glaciers and the polar ice sheets thinned and melted at their fringes. On average around the globe, sea level is estimated to have risen at rates exceed-ing 2.5 mm/year between about 8,000 and 6,000 years before present. These rates steadily decreased to less than 2.0 mm/year through about 4,000 years ago and stabilized at less than 0.4 mm/year through the late 1800s. Global sea level rise has accelerated again within the last 100 years, and now averages about 1 to 2 mm/year.133 See Chapter 12: Sea Level Rise for more thorough analysis of how sea level rise has already and will affect the U.S. coasts.

Figure 13.6: Predicted change in dissolved oxygen on the σθ = 26.5 (average depth of approximately 290 m) potential density surface, between the 1981–2000 and 2081–2100, based on the Community Earth System Models–Large En-semble Experiments (Figure source: redrawn from Long et al. 2016128).

mmol/m3

2001851701551401251109580655035205−200 −185 −170 −155 −140 −125−110 −95 −80 −65 −50 −35 −20 −5

Projected Change in Dissolved OxygenProjected Change in Dissolved Oxygen



13.5.2 Wet and Dry DepositionDust transported from continental desert re-gions to the marine environment deposits nu-trients such as iron, nitrogen, phosphorus, and trace metals that stimulate growth of phyto-plankton and increase marine productivity.134 U.S. continental and coastal regions experi-ence large dust deposition fluxes originating from the Saharan desert to the East and from Central Asia and China to the Northwest.135 Changes in drought frequency or intensity re-sulting from anthropogenically forced climate change, as well as other anthropogenic activ-ities such as agricultural practices and land-use changes may play an important role in the future viability and strength of these dust sources (e.g., Mulitza et al. 2010136).

Additionally, oxidized nitrogen, released during high-temperature combustion over land, and reduced nitrogen, released from intensive agriculture, are emitted in high pop-ulation areas in North America and are carried away and deposited through wet or dry depo-sition over coastal and open ocean ecosystems via local wind circulation. Wet deposition of pollutants produced in urban areas is known to play an important role in changes of eco-system structure in coastal and open ocean systems through intermediate changes in the biogeochemistry, for instance in dissolved oxygen or various forms of carbon.137

13.5.3 Primary ProductivityMarine phytoplankton represent about half of the global net primary production (NPP) (approximately 50 ± 28 GtC /year), fixing atmospheric CO2 into a bioavailable form for utilization by higher trophic levels (see also Ch. 2: Physical Drivers of Climate Change).138,

139 As such, NPP represents a critical compo-nent in the role of the oceans in climate feed-back. The effect of climate change on primary productivity varies across the coasts depend-ing on local conditions. For instance, nutrients

that stimulate phytoplankton growth are impacted by various climate conditions, such as increased stratification which limits the transport of nutrient-rich deep water to the surface, changes in circulation leading to vari-ability in dry and wet deposition of nutrients to coasts, and altered precipitation/evapora-tion which changes runoff of nutrients from coastal communities. The effect of the multiple physical factors on NPP is complex and leads to model uncertainties.140 There is consider-able variation in model projections for NPP, from estimated decreases or no changes, to the potential increases by 2100.141, 142, 143 Simula-tions from nine Earth system models projected total NPP in 2090 to decrease by 2%–16% and export production (that is, particulate flux to the deep ocean) to drop by 7%–18% as com-pared to 1990 (RCP8.5).142 More information on phytoplankton species response and asso-ciated ecosystem dynamics is needed as any reduction of NPP and the associated export production would have an impact on carbon cycling and marine ecosystems.

13.5.4 EstuariesEstuaries are critical ecosystems of biological, economic, and social importance in the United States. They are highly dynamic, influenced by the interactions of atmospheric, freshwater, terrestrial, oceanic, and benthic components. Of the 28 national estuarine research reserves in the United States and Puerto Rico, all are being impacted by climate change to varying levels.144 In particular, sea level rise, saltwa-ter intrusion, and the degree of freshwater discharge influence the forces and processes within these estuaries.145 Sea level rise and subsidence are leading to drowning of existing salt marshes and/or subsequent changes in the relative area of the marsh plain, if adaptive upslope movement is impeded due to urban-ization along shorelines. Several model sce-narios indicate a decline in salt marsh habitat quality and an accelerated degradation as the



rate of sea level rise increases in the latter half of the century.146, 147 The increase in sea level as well as alterations to oceanic and atmospheric circulation can result in extreme wave con-ditions and storm surges, impacting coastal communities.144 Additional climate change impacts to the physical and chemical estuarine processes include more extreme sea surface temperatures (higher highs and lower lows compared to the open ocean due to shallower depths and influence from land temperatures), changes in flow rates due to changes in pre-cipitation, and potentially greater extents of salinity intrusion.



TRACEABLE ACCOUNTSKey Finding 1The world’s oceans have absorbed about 93% of the excess heat caused by greenhouse gas warming since the mid-20th century, making them warmer and alter-ing global and regional climate feedbacks. Ocean heat content has increased at all depths since the 1960s and surface waters have warmed by about 1.3° ± 0.1°F (0.7° ± 0.08°C) per century globally since 1900 to 2016. Under a higher scenario, a global increase in average sea surface temperature of 4.9° ± 1.3°F (2.7° ± 0.7°C) by 2100 is projected, with even higher changes in some U.S. coastal regions. (Very high confidence)

Description of evidence baseThe key finding and supporting text summarizes the evidence documented in climate science literature, in-cluding Rhein et al. 2013.4 Oceanic warming has been documented in a variety of data sources, most nota-bly the World Ocean Circulation Experiment (WOCE) (http://www.nodc.noaa.gov/woce/wdiu/) and Argo databases (https://www.nodc.noaa.gov/argo/) and Extended Reconstructed Sea Surface Temperature (ERSST) v4 (https://www.ncdc.noaa.gov/data-access/marineocean-data/extended-reconstructed-sea-sur-face-temperature-ersst-v4). There is particular confi-dence in calculated warming for the time period since 1971 due to increased spatial and depth coverage and the level of agreement among independent SST ob-servations from satellites, surface drifters and ships, and independent studies using differing analyses, bias corrections, and data sources.6, 7, 11 Other observa-tions such as the increase in mean sea level rise (see Ch. 12: Sea Level Rise) and reduced Arctic/Antarctic ice sheets (see Ch. 11: Arctic Changes) further confirm the increase in thermal expansion. For the purpose of extending the selected time periods back from 1900 to 2016 and analyzing U.S. regional SSTs, the ERSST version 4 (ERSSTv4)19 is used. For the centennial time scale changes over 1900–2016, warming trends in all regions are statistically significant with the 95% con-fidence level. U.S. regional SST warming is similar be-tween calculations using ERSSTv4 in this report and those published by Belkin,148 suggesting confidence in these findings. The projected increase in SST is based

on evidence from the latest generation of Earth System Models (CMIP5).

Major uncertaintiesUncertainties in the magnitude of ocean warming stem from the disparate measurements of ocean tempera-ture over the last century. There is low uncertainty in warming trends of the upper ocean temperature from 0–700 m depth, whereas there is more uncertainty for deeper ocean depths of 700–2,000 m due to the short record of measurements from those areas. Data on warming trends at depths greater than 2,000 m are even more sparse. There are also uncertainties in the timing and reasons for particular decadal and interan-nual variations in ocean heat content and the contri-butions that different ocean basins play in the overall ocean heat uptake.

Summary sentence or paragraph that integrates the above informationThere is very high confidence in measurements that show increases in the ocean heat content and warm-ing of the ocean, based on the agreement of different methods. However, long-term data in total ocean heat uptake in the deep ocean are sparse leading to limited knowledge of the transport of heat between and with-in ocean basins.

Key Finding 2The potential slowing of the Atlantic Meridional Over-turning Circulation (AMOC; of which the Gulf Stream is one component)—as a result of increasing ocean heat content and freshwater driven buoyancy changes—could have dramatic climate feedbacks as the ocean absorbs less heat and CO2 from the atmosphere.51 This slowing would also affect the climates of North Amer-ica and Europe. Any slowing documented to date can-not be directly tied to anthropogenic forcing primarily due to lack of adequate observational data and to chal-lenges in modeling ocean circulation changes. Under a higher scenario (RCP8.5) in CMIP5 simulations, the AMOC weakens over the 21st century by 12% to 54% (low confidence).



Description of evidence baseInvestigations both through direct observations and models since 20134 have raised significant concerns about whether there is enough evidence to determine the existence of an overall slowdown in the AMOC. As a result, more robust international observational cam-paigns are underway currently to measure AMOC circu-lation. Direct observations have determined a statistical-ly significant slowdown at the 95% confidence level at 26°N (off Florida; see Baringer et al. 2016149) but model-ing studies constrained with observations cannot attri-bute this to anthropogenic forcing.39 The study29 which seemed to indicate broad-scale slowing has since been discounted due to its heavy reliance on sea surface tem-perature cooling as proxy for slowdown rather than ac-tual direct observations. Since Rhein et al. 2013,4 more observations have led to increased statistical confidence in the measurement of the AMOC. Current observation trends indicate the AMOC slowing down at the 95% con-fidence level at 26°N and 41°N but a more limited in situ estimate at 35°S, shows an increase in the AMOC.35, 149 There is no one collection spot for AMOC-related data, but the U.S. Climate Variability and Predictability Pro-gram (US CLIVAR) has a U.S. AMOC priority focus area and a webpage with relevant data sites (https://usclivar.org/amoc/amoc-time-series).

The IPCC 2013 WG1 projections indicate a high likeli-hood of AMOC slowdown in the next 100 years, how-ever overall understanding is limited by both a lack of direct observations (which is being remedied) and a lack of model skill to resolve deep ocean dynamics. As a result, this key finding was given an overall assessment of low confidence.

Major uncertaintiesAs noted, uncertainty about the overall trend of the AMOC is high given opposing trends in northern and southern ocean time series observations. Although earth system models do indicate a high likelihood of AMOC slowdown as a result of a warming, climate projections are subject to high uncertainty. This un-certainty stems from intermodel differences, internal variability that is different in each model, uncertainty in stratification changes, and most importantly uncer-

tainty in both future freshwater input at high latitudes as well as the strength of the subpolar gyre circulation.

Summary sentence or paragraph that integrates the above informationThe increased focus on direct measurements of the AMOC should lead to a better understanding of 1) how it is changing and its variability by region, and 2) whether those changes are attributable to climate driv-ers through both model improvements and incorpora-tion of those expanded observations into the models.

Key Finding 3The world’s oceans are currently absorbing more than a quarter of the CO2 emitted to the atmosphere annually from human activities, making them more acidic (very high confidence), with potential detrimental impacts to marine ecosystems. In particular, higher-latitude sys-tems typically have a lower buffering capacity against pH change, exhibiting seasonally corrosive conditions sooner than low-latitude systems. Acidification is re-gionally increasing along U.S. coastal systems as a re-sult of upwelling (for example, in the Pacific Northwest) (high confidence), changes in freshwater inputs (for example, in the Gulf of Maine) (medium confidence), and nutrient input (for example, in agricultural water-sheds and urbanized estuaries) (high confidence). The rate of acidification is unparalleled in at least the past 66 million years (medium confidence). Under the high-er scenario (RCP8.5), the global average surface ocean acidity is projected to increase by 100% to 150% (high confidence).

Description of evidence baseEvidence on the magnitude of the ocean sink is ob-tained from multiple biogeochemical and transport ocean models and two observation-based estimates from the 1990s for the uptake of the anthropogen-ic CO2. Estimates of the carbonate system (DIC and alkalinity) were based on multiple survey cruises in the global ocean in the 1990s (WOCE – now GO-SHIP, JGOFS). Coastal carbon and acidification surveys have been executed along the U.S. coastal large marine eco-system since at least 2007, documenting significantly elevated pCO2 and low pH conditions relative to oce-

https://usclivar.org/amoc/amoc-time-series

https://usclivar.org/amoc/amoc-time-series



anic waters. The data are available from the National Centers for Environmental Information (https://www.ncei.noaa.gov/). Other sources of biogeochemical bot-tle data can be found from HOT-DOGS ALOHA (http://hahana.soest.hawaii.edu/hot/hot-dogs) or CCHDO (https://cchdo.ucsd.edu/). Rates of change associated with the Palaeocene-Eocene Thermal Maximum (PETM, 56 million years ago) were derived using stable carbon and oxygen isotope records preserved in the sedimen-tary record from the New Jersey shelf using time se-ries analysis and carbon cycle–climate modelling. This evidence supports a carbon release during the onset of the PETM over no less than 4,000 years, yielding a maximum sustained carbon release rate of less than 1.1 GtC per year.86 The projected increase in global surface ocean acidity is based on evidence from ten of the lat-est generation earth system models which include six distinct biogeochemical models that were included in the latest IPCC AR5 2013.

Major uncertaintiesIn 2014 the ocean sink was 2.6 ± 0.5 GtC (9.5 GtCO2), equivalent to 26% of the total emissions attributed to fossil fuel use and land use changes.67 Estimates of the PETM ocean acidification event evidenced in the geo-logical record remains a matter of some debate within the community. Evidence for the 1.1 GtC per year cit-ed by Zeebe et al.,86 could be biased as a result of brief pulses of carbon input above average rates of emis-sions were they to transpire over timescales ≲ 40 years.

Summary sentence or paragraph that integrates the above informationThere is very high confidence in evidence that the oceans absorb about a quarter of the carbon dioxide emitted in the atmosphere and hence become more acidic. The magnitude of the ocean carbon sink is known at a high confidence level because it is estimated using a series of disparate data sources and analysis methods, while the magnitude of the interannual variability is based only on model studies. There is medium confidence that the current rate of climate acidification is unprecedented in the past 66 million years. There is also high confidence that oceanic pH will continue to decrease.

Key Finding 4Increasing sea surface temperatures, rising sea levels, and changing patterns of precipitation, winds, nutri-ents, and ocean circulation are contributing to over-all declining oxygen concentrations at intermediate depths in various ocean locations and in many coastal areas. Over the last half century, major oxygen loss-es have occurred in inland seas, estuaries, and in the coastal and open ocean (high confidence). Ocean ox-ygen levels are projected to decrease by as much as 3.5% under the higher scenario (RCP8.5) by 2100 rela-tive to preindustrial values (high confidence).

Description of evidence baseThe key finding and supporting text summarizes the evidence documented in climate science literature including Rhein et al. 2013,4 Bopp et al. 2013,24 and Schmidtko et al. 2017.150 Evidence arises from extensive global measurements of the WOCE after 1989 and in-dividual profiles before that.94 The first basin-wide dis-solved oxygen surveys were performed in the 1920s.150 The confidence level is based on globally integrated O2 distributions in a variety of ocean models. Although the global mean exhibits low interannual variability, re-gional contrasts are large.

Major uncertaintiesUncertainties (as estimated from the intermodel spread) in the global mean are moderate mainly be-cause ocean oxygen content exhibits low interannual variability when globally averaged. Uncertainties in long-term decreases of the global averaged oxygen concentration amount to 25% in the upper 1,000 m for the 1970–1992 period and 28% for the 1993–2003 period. Remaining uncertainties relate to regional vari-ability driven by mesoscale eddies and intrinsic climate variability such as ENSO.

Summary sentence or paragraph that integrates the above informationMajor ocean deoxygenation is taking place in bodies of water inland, at estuaries, and in the coastal and the open ocean (high confidence). Regionally, the phenom-enon is exacerbated by local changes in weather, ocean circulation, and continental inputs to the oceans.



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